L e
A Introduction
Nous avons vu dans le chapitre 3 que de nombreux modèles étaient proposés pour expliquer l’évolution du Junggar Occidental au cours du Paléozoïque (Buckman & Aitchison, 2004 ; Geng et al., 2009 ; Zhang et al., 2011a ). Le modèle de subduction de ride (Geng et al., 2009) est essentiellement basée sur les données géochimiques alors que le modèle d’accrétions/collisions successives (Buckman & Aitchison, 2004) s’appuie sur des contraintes lithologiques et géochronologiques. Ces modèles permettent d’expliquer la formation du complexe d’accrétion d’âge Paléozoïque supérieur et le magmatisme contemporain, mais la structure du complexe reste très mal connue. Sur la base de données structurales, Zhang et al. (2011a) proposent une double subduction au carbonifère, mais nous avons vu précédemment que cette interprétation était discutable. L’objectif de cette étude est d’apporter des observations de terrain sur la structure du complexe d’accrétion.
Parallèlement, les modèles cinématiques proposés en Asie Centrale montrent la présence d’une virgation au Kazakhstan pendant le Paléozoïque supérieur (Abrajevitch et al., 2007 ; 2008). La marge active associé à cette virgation est responsable de la disparition par subduction de l’Océan Junggar Balkash (Filippova et al., 2001). La virgation kazakh se connecte à l’Ouest Junggar et au nord Tianshan, si bien que l’évolution de ces deux régions est vraisemblablement due à la cinématique régionale affectant les Altaïdes Occidentales à la fin du Paléozoïque. Cette période est aussi marquée par la transition entre le régime compressif entraînant la subduction dévono-carbonifère et le régime transcurrent à l’origine des mouvements relatifs entre les blocs d’Asie Centrale (van der Voo et al., 2006 ; Wang et al., 2007). Ces mouvements sont principalement des rotations de blocs autour d’un axe vertical qui sont accommodés par de grands décrochements documentés dans l’Altaï (Laurent-Charvet et al., 2002 ; 2003 ; Buslov et al., 2004), dans le Tianshan (Yin & Nie, 1996 ; Laurent-Charvet et al., 2002 ; Wang et al., 2010) et au Kazakhstan. Le second objectif de cette étude est de caractériser, à partir de l’étude structurale d’un segment de l’Asie Centrale, la transition entre la subduction carbonifère et le régime décrochant permien. Ce travail permettra ainsi de replacer l’évolution fini-paléozoïque du Junggar Occidental au sein du cadre géodynamique de l’Asie Centrale.
B Article accepté à Gondwana Research: From oblique accretion to transpression in the evolution of the Altaid collage: new insights from West Junggar, northwestern China
Flavien Choulet1,*, Michel Faure1, Dominique Cluzel2,Yan Chen1,Wei Lin3, Bo Wang4 1
: Institut des sciences de la Terre d’Orléans, UMR 6113 - CNRS/Université d'Orléans 1A, rue de la Férollerie, 45071 Orléans CEDEX 2, France
2
: Pôle Pluridisciplinaire de la Matière et de l'Environnement, EA 3325, Université de la Nouvelle-Calédonie, Nouméa, Nouvelle-Calédonie, France
3
: State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, P.O. Box 9825, Beijing 100029, China.
4
: Department of Earth Sciences, Nanjing University, Nanjing, China *: flavien.choulet@univ-orleans.fr
Abstract
Along active margins, tectonic features that develop in response to plate convergence are strongly controlled by subduction zone geometry. In West Junggar, a segment of the giant Palaeozoic collage of Central Asia, the West Karamay Unit represents a Carboniferous accretionary complex composed of fore-arc sedimentary rocks and ophiolitic mélanges. The occurrence of quasi-synchronous upright folds and folds with vertical axes suggests that transpression plays a significant role in the tectonic evolution of the West Junggar. Latest Carboniferous (ca. 300 Ma) alkaline plutons postdate this early phase of folding, which was synchronous with accretion of the Carboniferous complex. The Permian Dalabute sinistral fault overprints Carboniferous ductile shearing and split the West Karamay Unit ca. 100 kilometres apart. Oblique convergence may have been provoked by the buckling of the Kazakh orocline and relative rotations between its segments. Depending upon the shape of the convergence zone, either upright folds and fold with vertical axes, or alternatively, strike-slip brittle faults developed in response to strain partitioning. Sinistral brittle faulting may account for the lateral imbrication of units in the West Junggar accretionary complex.
Keywords
B.1 Introduction
In contrast with strictly frontal convergence, which is rarely observed, examples of oblique subduction are widespread (Chamot-Rooke & Rabaute, 2007), and often generate strike-slip faults parallel to the upper plate boundary (Allen, 1965; Katili, 1970). The western North American Cordilleras, Andes, Taiwan, and Sumatra are the best examples of such an oblique convergent setting. Fitch (1972) was the first to link the tectonic structures in the upper plate to the oblique slip of the lower plate. Based on earthquake focal mechanisms in western Pacific, he proposed that the total decoupling of the oblique slip would result into a component of convergence normal to the trench and a shearing component parallel to the trench marked by transcurrent faulting. Beck (1983) improved this model by establishing the geometric and thermal constraints that favour decoupling of oblique convergence. Very oblique convergence, gently dipping subduction and thermal softening of the upper plate are the main conditions that favour the decoupling of oblique slip in a subduction zone.
Because total decoupling of oblique convergence is rarely achieved at sites of oceanic subduction, McCaffrey (1992) proposed a partial decoupling model, and demonstrated that margin geometry could influence the tectonic response of the upper plate. Therefore, oblique convergence along a concave or a convex subduction zone toward the ocean will be accommodated by transpression or transtension, respectively. The present curvature of the western Sunda and Aleoutian subduction zones (Ekström & Engdahl, 1989; McCaffrey, 1991) are good paradigms of oblique slip partitioning that may also be reproduced by analogical modelling (Chemenda et al., 2000). The rheology of the accretionary wedge also influences the geometric variability of the subduction zone (Platt, 1993b). Very oblique convergence would logically generate an intense slicing of the upper plate boundary (Martinez et al., 2002). Triple junctions and ridge subduction can also account for the initiation or reactivation of strike-slip faults in the overriding plate (Thorkelson, 1996; Roeske et al., 2003).
Lateral tectonic transport along the active margin is a direct consequence of decoupling (Coney et al., 1980; Beck, 1983; Jarrard, 1986); it is referred as “Sunda style” tectonics (Beck, 1983), and thousands of kilometres along-margin displacements have been evidenced in far-travelled allochthonous terranes of western North America (Beck, 1980; Coney et al., 1980). However, in most cases, terrane traveling is limited to a few tens of kilometres (Beck, 1986). This variability depends upon the age and obliquity of the
that may result in lateral terrane transport significantly contributes to lateral growth of the continental margin and, consequently, to a reorganisation of the continental crust pattern.
During the last decades, Mesozoic and Cenozoic cases of oblique subduction have been established in the Circum-Pacific area, (Karig et al., 1978; Engebretson et al., 1985; Kimura, 1986; Reutter et al., 1991; Beck, 1994; Kusky et al., 1997a, 1997b) by comparison with modern analogues (Malod et al., 1995; Lallemand et al., 1999; Goldfinger et al., 1996). In contrast, oblique subduction is rarely documented in older accretionary orogens (Henderson, 1987; Veevers, 2003). The purpose of this article is to report an example of Palaeozoic oblique convergence and to discuss its regional geodynamic controls.
The Altaids (Sengör et al., 1993; Sengör & Natal’in, 1996a) or Central Asian Orogenic Belt (CAOB; Mossakovsky et al., 1993; Windley et al., 2007) are a wide orogenic collage formed during the Palaeozoic as a result of the convergence of Siberia, Baltica, Tarim, and North China blocks (Fig. 6.B.1a). Because of post-Palaeozoic tectonics, the present structure exhibits a distorted pattern of accretionary complexes, magmatic arcs, and ribbon-like micro continents. Several conflicting models have been proposed for the Altaids (for a review see Windley et al., 2007 and Xiao et al., 2010). The Kipchak Arc model is
Figure 6.B.1: a) location of the Altaids including major cratons and orogenic belts of Eurasia. b) structural map of western Altaids, modified after Windley et al. (2007) and Charvet et al. (2007). The Devonian to Carboniferous Kazakh orocline lying on the pre-Devonian Kazakhstan microcontinent is the major structure of this region. The nature of the microcontinent in the core of the orocline, below the Junggar basin is still controversial, and a discussion on this topic is beyond the scope of this paper. Major faults are also represented. BOLE: Bole Block, CANTF: Chingiz-Alakol-North Tianshan, CKF: Central Kazakhstan Fault, DF: Dalabute Fault, IGSZ: Irtysh-Gornotsaev Shear Zone, MTF: Main Tianshan Fault, NNTL: Nalati-Nikolaiev Teconic Line, TTF: Talas-Fergana Fault.
characterized by a single long-lived subduction that was later shredded by strike-slip faults (Sengör et al., 1993; Sengör & Natal’in, 1996a). An archipelago model was alternatively proposed (Filippova et al., 2001; Xiao et al., 2008); it consists of accreted and laterally docked pairs of associated accretionary complexes and magmatic arcs. A remarkable feature of the Altaids is the presence of horseshoe-shaped belts, such as the Kazakh Orocline (Fig.
6.B.1b; Abrajevitch et al., 2008), or the Central Mongol Orocline (Yakubchuk et al., 2008).
These structures are intimately associated with lithosphere-scale strike-slip faults along which palaeomagnetic evidence document block rotations and displacements over thousands kilometres (van der Voo et al., 2006; Wang et al., 2007; Choulet et al., 2010); however, the link between oroclinal bending, transcurrent faulting and accretion remains poorly understood.
This study deals with the structural pattern of the Late Palaeozoic West Karamay accretionary complex, in order to document transcurrent tectonics and lateral docking. On the basis of new geochronological data and multi-scale structural analysis, we present the first evidence of an oblique convergent system in West Junggar. Considering the structural pattern of the Central Asian puzzle, we discuss the possible origin of oblique subduction, and the controls of regional geodynamics on the geometry of the convergent plate boundary.
B.2 Geological outline
B.2.i Central Asia
In the central part of the Altaids, a region that extends from central Kazakhstan to Xinjiang (northwestern China), three main geological domains are recognized (Fig. 6.B.1b). To the northeast, (1) the Altai range is composed by Early and Late Palaeozoic units that were accreted and docked to the Siberian margin and affected by high-grade metamorphism (Windley et al., 2002; Xiao et al., 2004). To the south, the convergence between the Tarim Block and several micro continents such as Yili and Central Tianshan formed the (2) Palaeozoic Tianshan Orogen (Charvet et al., 2007). The central and northwestern parts of Central Asia display a horseshoe shape that can be followed from North Tianshan to West Junggar around the Balkash Lake area (Fig. 6.B.1b). This megastructure is termed the (3) Kazakh Orocline (Zonenshain et al., 1990). In central Kazakhstan, the outer part of the orocline is made of micro continents and intra-oceanic arcs, which amalgamated during the Early Palaeozoic (Kröner et al., 2008). In the inner part of the orocline, the subduction of the Junggar Ocean below the Kazakhstan active margin generated Late Palaeozoic accretionary
complexes and magmatic arcs (Degtyarev, 1999; Wang et al., 2006; Windley et al., 2007). To the north of this domain (Fig. 6.B.1b), the Irtysh-Zaisan fold-and-thrust Belt results from the Late Carboniferous closure of the Ob-Zaisan Ocean that originally separated the Kazakh orocline and the south-western margin of Siberia (Buslov et al., 2004). The Permian-Early Triassic transcurrent tectonics that affected Central Asia (Allen et al., 1995; Laurent-Charvet et al., 2003), eventually dismembered the oroclinal system, displaced segments over more than 1000 km, and thus disorganised its original structure (Wang et al., 2007; Choulet et al., 2010).
B.2.ii West Junggar
West Junggar, a mountainous area located along the Kazakh border in northwestern China, forms the easternmost part of the Kazakh orocline (Fig. 6.B.1b). It is limited by two major strike-slip fault systems, the Irtysh-Gornotsaev sinistral shear zone to north and the Chingiz-Alakol-North Tianshan dextral shear zone to the south (Choulet et al., 2010; Fig.
6.B.1b). Permian displacements along these faults have been estimated from several hundreds
to more than one thousand kilometres (Wang et al., 2007; Choulet et al., 2010). These faults represent major tectonic boundaries between West Junggar, Altai, and Tianshan. Although detailed investigations are rare in West Junggar, several authors have recognized numerous stratigraphic and tectonic units (Feng et al., 1989; Buckman & Aitchison, 2004). The section below is a brief summary of the litho-stratigraphic units defined in Choulet et al. (unpublished results; Fig. 6.B.2a).
The Chingiz-Tarbagatay Unit in the central part of the West Junggar massif, is composed of Early Palaeozoic mélange, turbidite and magmatic arc rocks (Unit I in Fig.
6.B.2a; Feng et al., 1989). The Mayila and Tangbale Units are also formed by Early
Palaeozoic ophiolitic mélanges and turbidites (Units IVa and IVb in Fig. 6.B.2a; Buckman & Aitchison, 2004). Unconformable Middle Devonian conglomerate that overlie Ordovician and Silurian rocks argue for a Late Silurian event (XBGRM, 1965a). A-type Early Devonian granites intrude the Chingiz-Tarbagatay Unit and postdate the pre-Late Silurian accretion-subduction (Chen et al., 2010a). These units with still a poorly documented architecture represent the substratum of the Devonian-Carboniferous arcs (Units IIa and IIIa in Fig.
6.B.2a). At the end of the Middle Devonian, two new subduction zones developed. To the
north, the south-dipping subduction of the Ob-Zaisan Ocean generated the Sawuer arc and Erquis accretionary complex (Unit IIa and IIb in Fig. 6.B.2a; Windley et al., 2007; Shen et al., 2008; Zhou et al., 2008b; Chen et al., 2010a). To the south, the Barliek magmatic arc,
and the West Karamay accretionary complex are related to the northwest-dipping subduction of the Junggar Ocean (Units IIIa and IIIb in Fig 6.B.2a; Feng et al., 1989; Buckman &
Figure 6.B.2: a) map of West Junggar Mountains, showing the different tectonic units. Two pairs of Late Palaeozoic accretionary complexes and magmatic arcs overlie an Early Palaeozoic substratum, itself formed by arc magmatism and accretion. The location of samples described in this article is also presented. b) structural map of the West Karamay Unit. This unit is in fault contact with the surrounding Barliek, Mayila and Tangbale units. The West Karamay unit is an accretionary complex that comprises Early to Late Carboniferous sedimentary rocks (turbidite series and mass-flow greywacke deposits) and ophiolitic mélanges. The NE-SW trending Dalabute fault separates the unit in two parts. c) geological section across Barliek magmatic arc end West Karamay accretionary complex. P: Permian, Mz-Cz: Mesozoic and Cenozoic sedimentary rocks.
Aitchison, 2004; Chen et al., 2006; Xiao et al., 2008). This magmatic arc–subduction complex assemblage corresponds to the easternmost extension of the Kazakh orocline (Choulet et al., in press; Fig. 6.B.1b).
A particularity of West Junggar is the abundant and widespread Late Palaeozoic magmatism (Han et al., 2006; Fig 6.B.2a), which affected the entire Central Asia (Jahn et al., 2000a). Magmatic suites consist of A-type and I-type plutons, mafic dykes and volcanic rocks, emplaced between 320 Ma and 250 Ma (Chen & Jahn, 2004; Li et al., 2004a; Han et al., 2006; Xu et al., 2008b; Geng et al., 2009; Yin et al., 2010). A-type granitoids were generated either by a partial melting of the depleted-mantle reservoir (Han et al., 1999) or, alternatively, by a thermally induced melting of the Palaeozoic juvenile lower crust followed by in situ differentiation (Chen & Jahn, 2004; Su et al., 2006a); or both processes acting together (Chen & Arakawa, 2005; Geng et al., 2009). I-type granitoids stem from the melting of Early Palaeozoic juvenile crust (Chen & Jahn, 2004) or a depleted mantle reservoir (Zhou et al., 2008b). Dolerite and low-Mg diorite dykes dated between 283Ma and 241Ma (Qi, 1993; Li et al., 2004a; Xu et al., 2008b; Zhou et al., 2008a) have a depleted-mantle origin. All these rocks have been assigned a post-collisional setting.
In contrast, 320-300 Ma calk-alkaline rocks with adakitic affinities (Zhang et al., 2006; Geng et al., 2009; Tang et al., 2010) and high-Mg diorite dykes (Yin et al., 2010) were recently described and slab melting related to ridge subduction was proposed to account for their genesis. These new data led several authors to consider that subduction may have continued during Permian (Geng et al., 2010; Xiao et al., 2010), but this is not supported by field evidence. Actually, eruption of Permian lava flows (Tan et al., 2006) is closely associated with the accumulation of Permian coarse red sandstones and conglomerates, considered as a post-orogenic molasse (Feng et al., 1989; Allen et al., 1995; Jin & Li, 1999; Buckman & Aitchison, 2004; Fig. 6.B.2b). Undeformed Early Permian molasse postdates turbidite accumulation; therefore, subduction likely ended before the Early Permian.
All the Palaeozoic rocks of West Junggar have been affected by Permian post accretion transcurrent tectonics (Allen et al., 1995; Laurent-Charvet et al., 2003; Fig
6.B.1b). SW-NE trending faults, such as the Dalabute sinistral fault, affect Permian plutons
and generate cataclasite (Allen et al., 1995; Fig 6.B.2b), whilst ductile mylonite is never observed.
B.3 Age and nature of the West Karamay Unit
West Junggar Mountains are bounded to the east by the Junggar basin (Fig. 6.B.2b); in this area, low elevation and desert morphology expose discontinuous outcrops. From the bottom to the top, the Carboniferous Xibeikulasi, Baogoutu, and Tailegula formations have been classically recognized (XBGRM, 1966; 1978; Wu and Pan, 1991); however, similar lithologies and the lack of accurate stratigraphic evidence led several authors to reappraise this classification (Feng et al., 1989; Buckman & Aitchison, 2004; Choulet et al., unpublished results). In the following section, they will be collectively termed “West Karamay Unit” (Fig. 6.B.2b; Choulet et al., unpublished results). This unit consists of imbricate slices of turbidite, greywacke, and ophiolitic mélange (Feng et al., 1989), described thereafter (Fig. 6.B.2b).
B.3.i The turbidite series
In the West Karamay Unit, ca. 10 m-thick alternations of fine-grained grey siltstone and blackish mudstone are the predominant lithology (Feng et al., 1989; Li & Jin, 1989; Guo et al., 2002; Fig. 6.B.3a); in many places, the Permian magmatism and associated high heat flow transformed these rocks into hornfels (Choulet et al., unpublished results; Fig. 6.B.2a). In clastic rocks, quartz and clay are dominant, but many feldspar and lithic clasts are also preserved (Fig. 6.B.3b), and coarse-grained, greywacke contains numerous andesite clasts (Fig. 6.B.3c). Slumps, disrupted soft sandstone beds, Bouma sequences (Fig. 6.B.3d), and the coexistence of deep-water and shallow-water ichnofacies attest to the tectonic instability of the basin and resedimentation processes (Jin and Li, 1999).
Rare fossils of plants, corals and brachiopods do not provide a better age assignment than Carboniferous (XBGRM, 1966; Li & Jin, 1989; Wu & Pan, 1991). Recent U-Pb geochronological data on detrital zircons yield a maximum Late Carboniferous age (ca. 305 Ma) for turbidite deposition, which is close to the age of accretion (Choulet et al., unpublished results). The positive εHf values of these zircons argue for a juvenile origin consistent with an immature active margin (Choulet et al., unpublished results). The bedding (S0) is usually apparent in coarse-grained turbidites, but often undistinguishable from the slaty cleavage (S1) in black mudstone. Relationships between bedding and cleavage will be described and discussed later in this article. Turbidites often dip steeply, however upright folds hinges are rarely observed (Fig. 6.B.3e).
Figure 6.B.3: Photographs of turbidites and greywackes from the West Junggar sedimentary units. a: turbidites made of decimetre-scale alternation of medium to coarse-grained volcaniclastc sandstone (45.9808°N; 85.3093°E), b: microphotographs of lithic, feldspar and quartz clasts within turbidite sandstone (45.7233°N; 84.4516°E), c: fine-grained andesite clast frequently appearing within turbidites (45.7233°N; 84.4516°E), d: syn-sedimentary load casts structures of sandstone-siltstone beds in turbidites (45.8675°N; 84.6934°E), e: upright fold in the turbidites (45.8702°N; 85.2176°E), f: microphotograph of greywacke showing plagioclase, amphibole and pyroxene clasts within a clayey matrix (45.7214°N; 84.4593°E), g: andesite clast with well-expressed fluidal texture in greywacke mass-flow deposit (45.7214°N; 84.4593°E).
Table 6.B.1: La-ICPMS U-Pb detrital zircon data. *: Degree of discordance.
Analysis No. Th
(ppm) U
(ppm) Th/U
Ratios Ages (Ma)
Disc% * 207Pb / 206Pb ± 1σ 207Pb / 235U ± 1σ 206Pb / 238U ± 1σ 208Pb / 232Th ± 1σ 207Pb / 206Pb ± 1σ 207Pb / 235U ± 1σ 206Pb / 238U ± 1σ 208Pb / 232Th ± 1σ DJ155 (n=26) 1 DJ155 03 238.42 255.70 0.93 0.0526 0.0005 0.3674 0.0079 0.0507 0.0011 0.0138 0.0003 312 22 318 6 319 7 276 6 -2.14 2 DJ155 05 76.35 89.53 0.85 0.0533 0.0014 0.3927 0.0202 0.0534 0.0017 0.0148 0.0007 342 57 336 15 335 10 297 15 2.08 3 DJ155 06 51.66 77.20 0.67 0.0531 0.0005 0.3920 0.0074 0.0535 0.0012 0.0156 0.0003 334 22 336 5 336 7 313 7 -0.52 4 DJ155 07 406.58 393.41 1.03 0.0532 0.0005 0.3956 0.0070 0.0539 0.0012 0.0160 0.0003 337 22 338 5 339 7 322 6 -0.53 5 DJ155 08 94.31 120.14 0.79 0.0528 0.0005 0.3717 0.0063 0.0510 0.0011 0.0137 0.0003 322 22 321 5 321 7 275 5 0.21 6 DJ155 09 84.89 102.06 0.83 0.0525 0.0005 0.3550 0.0070 0.0490 0.0011 0.0124 0.0003 308 21 309 5 309 7 249 6 -0.31 7 DJ155 10 52.99 74.93 0.71 0.0533 0.0007 0.3901 0.0113 0.0531 0.0013 0.0167 0.0005 339 28 334 8 334 8 335 10 1.73 8 DJ155 11 22.70 44.06 0.52 0.0531 0.0018 0.3821 0.0235 0.0522 0.0019 0.0172 0.0012 334 73 329 17 328 11 344 23 1.86 9 DJ155 12 359.05 362.73 0.99 0.0529 0.0006 0.3702 0.0050 0.0508 0.0011 0.0142 0.0002 325 23 320 4 319 6 284 5 1.85 10 DJ155 13 92.75 134.05 0.69 0.0530 0.0005 0.3778 0.0058 0.0517 0.0011 0.0145 0.0003 328 22 325 4 325 7 291 5 0.75 11 DJ155 15 28.56 56.07 0.51 0.0531 0.0006 0.3895 0.0096 0.0532 0.0012 0.0175 0.0005 335 24 334 7 334 7 351 10 0.26 12 DJ155 18 280.90 280.11 1.00 0.0529 0.0005 0.3758 0.0055 0.0515 0.0011 0.0150 0.0003 324 24 324 4 324 7 301 5 -0.12 13 DJ155 19 91.39 162.90 0.56 0.0528 0.0010 0.3596 0.0141 0.0494 0.0014 0.0134 0.0006 321 44 312 11 311 8 269 13 3.15 14 DJ155 21 99.04 122.76 0.81 0.0524 0.0005 0.3484 0.0067 0.0482 0.0010 0.0138 0.0003 305 24 303 5 303 6 277 6 0.44 15 DJ155 22 65.25 408.37 0.16 0.0548 0.0006 0.4894 0.0070 0.0648 0.0013 0.0197 0.0005 404 24 404 5 404 8 395 11 0 16 DJ155 24 10.73 26.74 0.40 0.0542 0.0008 0.4349 0.0138 0.0582 0.0015 0.0184 0.0008 378 31 367 10 365 9 368 15 3.49 17 DJ155 25 93.09 172.86 0.54 0.0529 0.0005 0.3783 0.0073 0.0519 0.0011 0.0139 0.0003 322 21 326 5 326 7 279 7 -1.25 18 DJ155 27 365.07 403.30 0.91 0.0527 0.0006 0.3665 0.0044 0.0504 0.0010 0.0142 0.0002 318 26 317 3 317 6 286 4 0.22 19 DJ155 30 166.33 158.89 1.05 0.0526 0.0005 0.3566 0.0057 0.0492 0.0010 0.0135 0.0002 309 21 310 4 310 6 272 5 -0.08 20 DJ155 32 25.58 47.25 0.54 0.0532 0.0006 0.3840 0.0098 0.0523 0.0012 0.0160 0.0005 339 25 330 7 329 7 322 10 2.98 21 DJ155 33 49.81 80.35 0.62 0.0527 0.0006 0.3599 0.0099 0.0495 0.0012 0.0136 0.0004 317 28 312 7 311 7 273 9 1.84 22 DJ155 34 123.65 133.63 0.93 0.0532 0.0005 0.3939 0.0085 0.0536 0.0012 0.0146 0.0003 339 22 337 6 337 7 292 7 0.64 23 DJ155 35 52.14 70.81 0.74 0.0528 0.0006 0.3683 0.0091 0.0506 0.0012 0.0124 0.0004 319 23 318 7 318 7 248 7 0.35 24 DJ155 36 77.59 100.61 0.77 0.0526 0.0006 0.3548 0.0085 0.0489 0.0011 0.0137 0.0004 312 25 308 6 308 7 275 7 1.38 25 DJ155 38 141.05 179.12 0.79 0.0527 0.0005 0.3673 0.0065 0.0506 0.0011 0.0139 0.0003 315 20 318 5 318 7 278 6 -0.97 26 DJ155 39 277.83 260.72 1.07 0.0527 0.0005 0.3705 0.0076 0.0510 0.0011 0.0142 0.0003 317 22 320 6 320 7 284 6 -1.05 C HA P IT R E VI : L’ EVO LU TI O N G EO D YN AM IQ UE DE L ’O UE S T J UN GGA R A U P ALEO ZO ÏQ U E S U PERI EU R M ECAN IS M ES ET EVO L U T ION DE S C HA INE S D ’ ACCRET IO N EN A SI E C EN T RAL E
B.3.ii The graywacke mass flows
Mass flows are lenses without obvious internal structure intercalated within turbidite series (Wu & Pan, 1991; Guo et al., 2002). These discharges of sand-sized volcanic materials can reach tens of metres in thickness. Despite highly variable geometry of the mass flow itself, the greywacke is very homogenous and occasionally well sorted (Guo et al., 2002); quartz, zoned feldspar and volcanic-rock clasts are dispersed in a matrix of fine-grained quartz and clay (Fig. 6.B.3f). Rock fragments are usually dark andesite, occasionally exposing fluidal texture (Fig. 6.B.3g), consistent with a volcanic-arc origin for these volcaniclastic rocks. Greywackes are lithologically identical to the volcaniclastic sandstone beds of turbidite sequences. Some of these rocks were previously referred to as volcanic tuffs (Wu & Pan, 1991; Buckman & Aitchison, 2004), however, the clayey matrix and rounded clasts clearly rule out a pyroclastic origin for this rocks,